Biogeosciences, 12, 7519–7533, 2015

www.biogeosciences.net/12/7519/2015/

doi:10.5194/bg-12-7519-2015

© Author(s) 2015. CC Attribution 3.0 License.

Coastal upwelling off Peru and Mauritania inferred from helium

isotope disequilibrium

R. Steinfeldt1, J. Sültenfuß1, M. Dengler2, T. Fischer2, and M. Rhein1

1Institute of Environmental Physics, University of Bremen, Bremen, Germany2GEOMAR, Helmholtz Centre for Ocean Research, Kiel, Germany

Correspondence to: R. Steinfeldt (rsteinf@physik.uni-bremen.de)

Received: 22 May 2015 – Published in Biogeosciences Discuss.: 14 July 2015

Revised: 17 November 2015 – Accepted: 2 December 2015 – Published: 21 December 2015

Abstract. Upwelling is an important process, bringing gases

and nutrients into the ocean mixed layer. The upwelling ve-

locities, however, are too small to be measured directly. Here

we use the surface disequilibrium of the 3He / 4He ratio mea-

sured in two coastal upwelling regions off Peru in the Pa-

cific ocean and off Mauritania in the Atlantic ocean to cal-

culate the regional distribution of vertical velocities. To also

account for the fluxes by diapycnal mixing, microstructure-

based observations of the vertical diffusivity have been per-

formed on all four cruises analysed in this study. The up-

welling velocities in the coastal regions vary between 1.1±

0.3× 10−5 and 2.8± 1.5× 10−5 m s−1 for all cruises. Ver-

tical velocities are also inferred from the divergence of the

wind-driven Ekman transport. In the coastal regimes, both

methods agree within the error range. Further offshore, the

helium-derived vertical velocity still reaches 1× 10−5 m s−1,

whereas the wind-driven upwelling from Ekman suction is

smaller by up to 1 order of magnitude. One reason for this

difference is ascribed to eddy-induced upwelling. Both ad-

vective and diffusive nutrient fluxes into the mixed layer

are calculated based on the helium-derived vertical veloci-

ties and the vertical diffusivities. The advective part of these

fluxes makes up at about 50 % of the total. The nutrient

flux into the mixed layer in the coastal upwelling regimes

is equivalent to a net community production (NCP) of

1.3± 0.3 g C m2 d−1 off Peru and 1.6–2.1± 0.5 g C m2 d−1

off Mauritania.

1 Introduction

Eastern boundary upwelling systems (EBUS), such as the

Canary, California, Humboldt, and Benguela Currents be-

long to the most productive marine ecosystems, e.g. Fréon

et al. (2009). The upwelling is caused by the wind-driven

surface circulation. Alongshore trade winds drive an offshore

Ekman flux, which leads to a horizontal flow divergence

at the coast and as a consequence upwelling of cold and

nutrient-rich subsurface water. They also transport climate-

relevant trace gases such as N2O from the ocean’s inte-

rior into the mixed layer and ultimately into the atmosphere

(Kock et al., 2012).

Crucial to quantify the role of upwelling for nutrient and

tracer budgets is the vertical velocity, by which substances

and gases are transferred from the subsurface into the mixed

layer. It is, however, much too small to be measured directly.

In previous studies upwelling has been derived from the wind

field via Ekman theory (McClain, 1993; Hagen, 2001), the

divergence of the horizontal velocity fields inferred from cur-

rent metres, drifters, and shipboard measurements (Wein-

gartner and Weisberg, 1991; Gouriou and Riverdin, 1992)

and geochemical tracers such as 14C (Broecker et al., 1978;

Quay et al., 1983; Toggweiler et al., 1991), pCO2 and AOU

(Wanninkhof et al., 1995), CFC-12 and SF6 (Tanhua and Liu,

2015) as well as the isotopes 7Be (Kadko and Johns, 2011;

Haskell II et al., 2015) and 3He (Klein and Rhein, 2004;

Rhein et al., 2010).

The helium method (Klein and Rhein, 2004; Rhein et al.,

2010) exploits the excess of the helium isotope 3He in

the upwelled waters to determine the vertical velocity. The

oceanic source for 3He is hydrothermal venting, mainly near

Published by Copernicus Publications on behalf of the European Geosciences Union.

7520 R. Steinfeldt et al.: Coastal upwelling

mid-ocean ridges, where primordial 3He is emitted (Lupton,

1983). The 3He enriched waters eventually reach the mixed

layer, e.g. by upwelling, where the excess 3He is outgassing

from the ocean.

Another process resulting in a net flux of properties from

the ocean interior into the mixed layer is diapycnal mixing.

Together with the helium measurements microstructure pro-

filing has been performed at a large number of stations, and

the diapycnal diffusivity has been inferred from the data off

Mauritania (Schafstall et al., 2010). These authors found ele-

vated dissipation rates of turbulent kinetic energy particularly

at the continental slope close to the shelf break.

Here we investigate (i) the coastal region off Peru from

5 to 16◦ S that is part of the Humboldt Current upwelling

system and (ii) the southern part of the Canary system off

Mauritania between 20 and 16◦ N. Both the Peruvian and the

Mauritanian upwelling regions have in common that they are

adjacent to oxygen minimum zones (Karstensen et al., 2008).

This oxygen minimum is most pronounced in the Humboldt

current south of 10◦ S, where denitrification within the up-

welling water occurs. By this process also N2O is produced,

whereas in the Mauritanian upwelling the main production

pathway for this substance is nitrification (Kock et al., 2012).

In contrast to the study of Rhein et al. (2010) at the equator,

in the upwelling region off Peru and Mauritania the Ekman

theory can be applied to infer the vertical velocity. After pre-

senting the data and methods, the Ekman and helium-derived

vertical velocities in the upwelling regions are compared.

Therefore, the cases for the open ocean (offshore region) and

the case of a lateral boundary (for the coastal region) are

considered. Also the role of eddies related to vertical veloc-

ities (eddy pumping) is investigated (McGillicuddy Jr. et al.,

2007). Then the contribution of the upwelling and diapycnal

mixing to the nutrient fluxes into the mixed layer are calcu-

lated.

2 Data

From the Peruvian upwelling area, about 300 helium sam-

ples have been taken at 62 stations during the cruise Me-

teor M91 in December 2012. In this region, upwelling occurs

throughout the year, with medium offshore transport in bo-

real autumn–early winter (Carr and Kearns, 2003). Helium

measurements in the upwelling region west of Mauritania

have been performed on three cruises (Fig. 1) The first cruise,

M68/3 on RV Meteor, was conducted during boreal summer

2006 (July–August), but was mainly located north of 18◦ N.

During the main upwelling season in boreal winter, two

cruises with helium data are available: P347 with the German

RV Poseidon took place in January 2007 and was restricted

to the near coastal region, whereas on ATA3 (French ves-

sel L’Aatalante, February 2008), a larger area was sampled

but with less spatial resolution. Altogether, about 500 helium

samples have been taken at 101 stations. All above men-

tioned cruises were part of the German research program

“Surface Ocean Processes in the Anthropocene (SOPRAN)”.

CTD−O2 profiles were collected using a Seabird SBE 911

system attached to a carousel water sampler with 10 L Niskin

bottles. Water subsamples from the Niskin bottles were used

for the analysis of biogeochemical properties (nutrients, he-

lium isotopes). From the biogeochemical parameters only

PO4 and helium isotopes are used in this study. The up-

welling areas off Peru and off Mauritania are located within

oxygen minimum zones (OMZ), and the low oxygen con-

centration due to the remineralization of organic matter is

correlated with high nutrient values. Off Peru, oxygen con-

centrations in the OMZ are so low that denitrification oc-

curs at some places. Therefore we consider here only phos-

phate fluxes and not nitrate to avoid having to deal with the

influence of the OMZ. On the four cruises, phosphate was

measured with different autoanalysers, the precision is about

0.02 µmolkg−1.

The isotopes 3He and 4He were analysed with the Bremen

high-resolution static mass spectrometer (Sültenfuß et al.,

2009). A very high resolution is necessary to distinguish be-

tween the mass-3 hydrogen species 1H2H (HD) and 3He. In

this study, the isotopic ratio 3He / 4He will be used as a tracer

for upwelled waters, which is expressed as δ3He [%], i.e. the

relative deviation from the atmospheric ratio:

δ3He[%] =

(3He/4He

)water−

(3He/4He

)air(

3He/4He)

air

· 100.

The measurement precision for the 3He / 4He ratio is in gen-

eral ±0.4 % or better (Sültenfuß et al., 2009). This value is

confirmed by the standard deviation of repeat samples taken

mainly on cruises M91 and P347.

Figure 1 shows the locations and δ3He in the mixed layer

for the M91 cruise in the Peruvian upwelling region and the

cruises M68/3, P347, and ATA3 off Mauritania. δ3He values

fall in different ranges for the two oceans. This is the result of

the difference of the helium-3 concentration in the subsurface

waters of the two upwelling regions (see Fig. 2). This differ-

ence can even be seen in the mixed layer, thus indicating the

entrainment of water into the mixed layer from below.

In order to distinguish between advective and diffusive

vertical 3He fluxes into the mixed layer, knowledge of the

diapycnal diffusivity is an important factor. This quantity

is determined from microstructure shear data that were col-

lected on all cruises using different tethered microstructure

profilers (MSS90L and MSS90D). Both instrument types are

equipped with two airfoil shear sensors, a fast temperature

sensor (FP07), an acceleration sensor, tilt sensors, and stan-

dard CTD sensors. For a detailed description see Prandke and

Stips (1998). From the small-scale velocity fluctuations mea-

sured by the MSS instruments, dissipation rates of turbulent

kinetic energy ε are derived by integrating shear wavenumber

spectra assuming isotropic turbulence. Processing details are

Biogeosciences, 12, 7519–7533, 2015 www.biogeosciences.net/12/7519/2015/

R. Steinfeldt et al.: Coastal upwelling 7521

M68/3 Jul 2006

26oW 24oW 22oW 20oW 18oW 16oW

16oN

18oN

20oN

22oN <−1.6%−1.6%− −1.4%−1.4% − −1.2%−1.2% − −1.0%

M91 Dec 2012

85oW 80oW 75oW 70oW 20oS

15oS

10oS

5oS

0o <−1.6%−1.6%− −1%−1% − −0.5%−0.5% − 0%0% − 0.5%0.5% − 1%>1%

ATA3 Feb 2008

20oW 18oW 16oW

16oN

18oN

20oN

<−1.6%−1.6%− −1.4%−1.4% − −1.2%−1.2% − −1.0%

P347 Jan 2007

18oW 30’ 17oW 30’ 16oW 30’ 15oW 17oN

30’

18oN

30’

19oN <−1.6%−1.6%− −1.4%−1.4% − −1.2%−1.2% − −1.0%−1.0% − −0.8%>−0.8%

(b)

(c)

(a)

(d)

Figure 1. δ3He [%] in the mixed layer for cruise M91 off Peru (a) and for cruises M68/3, P347 and ATA3 off Mauritania (b–d). Isobaths are

drawn every 1000 m, and the area of the “coastal region” is shaded grey (for details, see text). Note the different colour scale for the cruises

from the Peruvian (a) and Mauritanian (b–d) region. The error of the δ3He values in the mixed layer is 0.2 %.

described in Schafstall et al. (2010), where the microstruc-

ture shear data from the Mauritanian area are presented. Af-

ter applying corrections for unresolved spectral ranges and

loss of variance due to the finite sensor tip finally the diapyc-

nal diffusivity Kρ is inferred via the Osborn (Osborn, 1980)

relationship:

Kρ = 0ε

N2.

N denotes the local buoyancy frequency, and 0 the mixing

efficiency, which is set to a constant value of 0.2 (Oakey,

1982).

In addition, remote-sensing data of wind speed, primary

production and sea surface height are used in this study.

These data are available via the internet. The wind speed U10

is taken from the daily gridded ASCAT (ftp://ftp.ifremer.fr/

ifremer/cersat/products/gridded/mwf-ascat/data/daily) wind

product (for cruiuse M91 in 2012) and the older QuikSCAT

product (ftp://ftp.ifremer.fr/ifremer/cersat/products/gridded/

mwf-quikscat/data/daily, for the Mauritanian cruises in

2006–2008). For primary productivity, the 8 day MODIS

based estimates from http://www.science.oregonstate.edu/

ocean.productivity/index.php were used. The algorithm for

computing primary production is based on Behrenfeld

and Falkoweski (1997). Also used in this study are sea

level anomalies from the Aviso product (http://www.aviso.

oceanobs.com/duacs/).

3 Methods

In order to compute the upwelling velocity from the 3He dis-

equilibrium, the same box model as in Rhein et al. (2010) and

Klein and Rhein (2004) is applied. This model only has a ver-

tical dimension with two boxes. The upper box 1 represents

the mixed layer, where gas exchange with the atmosphere

takes place. The lower box 2 exchanges properties with box

1 by advective and diffusive vertical fluxes. This box thus

supplies the upwelling water enriched in 3He. In steady state,

the upward advective and diffusive fluxes of 3He from box 2

into the mixed layer (box 1) are compensated by outgassing

of 3He from box 1 into the atmosphere. This steady-state as-

sumption results in the following equation for inferring the

vertical velocity w:

0= Fg−Kv

dC

dz+w(C2−C1). (1)

The same 1-D box model has been used in (Tanhua and

Liu, 2015) for CFC-12 and SF6. Kadko and Johns (2011)

and Haskell II et al. (2015) use a modified version for the

inventory of 7Be. All these models neglect horizontal advec-

tion, i.e it is assumed that changes of the respective tracer in

the mixed layer is dominated by vertical processes. Another

interpretation is that Eq. (1) is valid in a Lagrangian coordi-

nate system moving with the surface patch (box 1), as long as

www.biogeosciences.net/12/7519/2015/ Biogeosciences, 12, 7519–7533, 2015

7522 R. Steinfeldt et al.: Coastal upwelling

the lower box 2 moves with a similar velocity and horizontal

mixing is small.

Applying the box model, the values of the gas exchange

rate Fg, the vertical diffusivity Kv (which is assumed to be

equal to the diapycnal diffusivity Kρ) and the vertical δ3He

gradient below the mixed layer dC/dz and the δ3He ratio C1

and C2 in boxes 1 and 2, have to be determined for each pro-

file. We mainly follow the procedure in Rhein et al. (2010),

but with some differences in detail.

The gas exchange rate

Fg = vg1C, (2)

is given by the gas transfer velocity vg and the helium-3 dis-

equilibrium in the mixed layer 1C = Ceq−C1. The equilib-

rium δ3He ratio Ceq is −1.6 %. The gas transfer velocity vg

has been calculated using the relationship given by Nightin-

gale et al. (2000):

vg = (3)

0.01/3600 · (0.222U210+ 0.333U3

10) · (Sc/600)−0.5

The gridded wind speed U10 is interpolated on the locations

of the helium measurements, and the daily values are aver-

aged over a time period of n days in advance of the sampling

date. n depends on the timescale of the gas exchange, i.e.

the mixed-layer depth and the gas transfer velocity itself. n

has been calculated as mean value for each cruise and varies

between 4 (cruise M68/3) and 8 (cruise P347) days.

The mixed-layer value C1 for δ3He (also for nutrients) is

calculated as the mean of all measurements within the mixed

layer at each station. For bottle data as helium and nutrients,

these are typically one or two data points per profile. The

mixed-layer depth is determined according to Levitus (1982)

with a density threshold of1σθ = 0.125 kgm−3. In addition,

a visual examination of each profile has been applied to avoid

an erroneous allocation of bottle data to the mixed layer. This

could be caused by the fact that the mixed-layer depth is de-

termined from CTD data during the downcast, whereas the

Niskin bottles are closed during the upcast.

In order to estimate the δ3He ratio of box 2 C2, the verti-

cal mean values of δ3He for each profile are computed over

an interval of 5 and 25 m below the mixed layer. The rel-

atively sparse bottle data (helium and nutrients) are verti-

cally interpolated onto 1 m intervals using a piecewise cu-

bic Hermite polynomial interpolation scheme that preserves

the shape of the data as in Tanhua et al. (2010), and then the

vertical mean value is calculated. The depth range from 5

to 25 m below the mixed layer is much smaller than in Rhein

et al. (2010), but comparable to those used in Tanhua and Liu

(2015) applying the same box model and in Schafstall et al.

(2010) and Kock et al. (2012) for calculating diffusive fluxes

of nitrate and N2O into the mixed layer. Directly at the base

of the mixed layer, vertical mixing might dominate (Kadko

and Johns, 2011), but the large concentration gradient of 3He

(and also nutrients) cannot be resolved by the coarse resolu-

tion of the bottle data. We thus determine the diffusive and

advective helium-3 flux from the data in box 2 and assume

that these fluxes are continuous into the mixed layer, i.e. no

flux divergence or convergence occurs in the “gap” of 5 m

between the boxes 1 and 2. In Rhein et al. (2010), C2 was

calculated as regional mean over several profiles. This is not

done here, as the regions adjacent to the coast shows a large

variability in C2 (see Fig. 2g and h), so the original values

for each helium-3 profile are retained.

Turbulent fluxes of 3He into the mixed layer are estimated

from the diffusion coefficient based on the microstructure

shear data and the vertical δ3He gradient. BothKv and dC/dz

in Eq. (1) are averaged over the same range 5–25 m below

the mixed layer as the 3He values for calculating C2. At

some profiles, only helium, but no microstructure data are

available. In these cases, the regional mean value of Kv over

coastal and offshore regions respectively has been used.

4 Property distribution in the coastal upwelling areas

The distribution of temperature, phosphate and helium-3 is

shown in Fig. 2 for a section along 8◦ S in the Peruvian up-

welling, the long 18◦ N section off Mauritania from cruise

M68/3 and a short section also along 18◦ N from cruise P347.

The ideal case of coastal upwelling is represented by the sec-

tion off Peru (Fig. 2, left column): the isolines of all prop-

erties and also the isopycnal charecterizing the central water

(σθ = 26.0 kgm−3) are lifted up towards the coast due to up-

welling, and the mixed layer becomes shallower. A similar

general feature can be observed along the 18◦ N section for

the cruise P347 (Fig. 2, right column), at least for the up-

per 50 m of the section and the isopycnal σθ = 26.0 kgm−3.

Near the coast below ≈ 50 m depth, the isolines are declin-

ing downward towards the coast for both cruises from the

Mauritanian upwelling.

The 18◦ N section of cruise M68/3 (Fig. 2, middle column)

has the largest offshore extension of all sections from the two

upwelling regions. The isolines and also the bottom of the

mixed layer show a conspicuous uplift towards the east be-

tween 24 and 20◦W. This is the location of the Canary Cur-

rent advecting water from the upwelling system further north,

which can be seen from the relatively low water temperature

in the mixed layer. The water within the coastal upwelling

region is fed from the south by South Atlantic Central Wa-

ter (SACW) (Hagen, 2001) and is more enriched in nutrients

and helium-3 and less saline (i.e. colder along isopycnals)

than the North Atlantic Central Water, which can be found

west of the Canary Current in the interior of the northern

subtropical gyre. The Central Water off Peru is clearly more

enriched in phosphate and helium-3 than off Mauritania.

In all three sections, δ3He values in the mixed layer are in

general larger than the equilibrium of −1.6 %. This is espe-

cially so in the Pacific due to the high helium-3 content of

Biogeosciences, 12, 7519–7533, 2015 www.biogeosciences.net/12/7519/2015/

R. Steinfeldt et al.: Coastal upwelling 7523

Pre

ssur

e [d

bar]

0

25

50

75

100

125

150

17.4°W 17°W 16.6°W

−2

0

2C1C2

δ3He[%]

−2−1.6−1.4−1.2−1−0.500.511.522.533.54

PO4[μmol/kg]

00.20.40.60.811.21.41.61.822.22.42.83.2

P347 18°Nθ

[°C]

131415161718192021222324252627

Pre

ssur

e [d

bar]

0

25

50

75

100

125

150

M68/3 18°N

δ3 He

[%]

80.8°W 80.4°W 80°W−2

0246

C1C2

26°W 24°W 22°W 20°W 18°W

−2

0

2C1C2

Pre

ssur

e [d

bar]

0

25

50

75

100

125

150

M91 8°S

(a) (b) (c)

(d) (e) (f)

(g) (h) (i)

(j) (k) (l)

Figure 2. Sections of potential temperature (a–c), phosphate (d–f), helium 3 (g–i) for one section from cruise M91 off Peru (along 8◦ S),

M68/3 (along 18◦ N) and P347 (also along 18◦ N). The grey line denotes the base of the mixed layer and the white line the isopycnal

σθ = 26.0 kgm−3. (j–l): mean δ3He in box 1 (mixed layer) and box 2 (5–25 m below the mixed layer) along the sections.

the upwelled waters. Offshore in the Atlantic, on some lo-

cations, equilibrium values are found, for instance west of

24◦W during cruise M68/3 (Fig. 2h).

The lowest row in Fig. 2 shows the δ3He values of box

1 (C1) (the mixed layer) and box 2 (C2) (5–25 m below the

mixed layer), which are used in Eq. (1) to infer the upwelling

velocities. For the 8◦ S section of cruise M91 off Peru, both

(C1) and (C2) are increasing landward, as would be expected

for enhanced upwelling at the coast. The difference between

(C1) and (C2) decreases in onshore direction from about 2 to

1 % in δ3He, indicating an enhanced exchange between both

layers by upwelling and/or diapycnal mixing. In the western

subtropical gyre area of the 18◦ N of cruise M68/3, the δ3He

values in box 1 and 2 are almost identical, as the NACW

is depleted in helium-3. Further east, where SACW is dom-

inating below the mixed layer, the difference between (C1)

and (C2) is larger but decreases towards the coast, as for

the section off Peru. Also for the much shorter 18◦ N sec-

tion of cruise P347 the upwelling and/or mixing near the

coast lead to a relatively small difference between the δ3He

values in box 1 and 2. The almost identical values of (C1)

and (C2) at 17◦W might result from a data gap below the

mixed layer (see the location of helium samples in Fig. 2i).

Note that some stations along this line have been repeated

within a few days, so for some locations two mean values

in box 1 and box 2 exist. For the Mauritanian upwelling,

δ3He within box 1 reaches the equilibrium value of−1.6 % at

about 17◦W for both cruises. However, during M68/3 further

west at some stations an oversaturated mixed layer concen-

tration of helium-3 has been observed, indicating upwelling

also at those offshore locations.

5 Results and discussion

5.1 Helium-derived upwelling velocities

The upwelling velocities calculated according to Eq. (1) are

shown in Fig. 3 for all four cruises. Negative values of w are

set to zero, so possible downwelling is not considered here.

Small differencesC1−C2 in the denominator of Eq. (1) result

in large upwelling velocities. In order to guarantee that such

high vertical velocities are not the consequence of uncertain-

ties in the helium measurement, values for w from Eq. (1)

are discarded if the absolute value |C1−C2| is smaller than

the quadratic sum of the uncertainties of C1 and C2, both

numbers are assumed to be 0.2 % in δ3He units (see next

subsection and Table 1).

Overall, at about 60 % of the stations with δ3He mea-

surements in the mixed-layer upwelling occurs (26 out of

49 stations for M91 and 43 out of 74 for the Mauritanian

cruises). The resulting vertical velocities are of the order of

www.biogeosciences.net/12/7519/2015/ Biogeosciences, 12, 7519–7533, 2015

7524 R. Steinfeldt et al.: Coastal upwelling

Table 1. Error estimation.

δ3He box 1 δ3He box 2 Vert. mixing Piston velocity vg Total

Uncertainty ±0.2 % ±0.2 % factor of 2 30 %

Error of wHe Peru ±22 % ±17 % ±42 % ±46 % ±68 %

Error of wHe Maur. ±42 % ±26 % ±64 % ±59 % ±100 %

ATA3 Feb 2008

20oW 18oW 16oW

16oN

18oN

20oN P347 Jan 2007

18oW 30’ 17oW 30’ 16oW 30’ 15oW 17oN

30’

18oN

30’

19oN 00−1 × 10−5m/s

1−2 × 10−5m/s

2−3 × 10−5m/s

3−5 × 10−5m/s

>5 × 10−5m/s

M91 Dec 2012

85oW 80oW 75oW 70oW 20oS

15oS

10oS

5oS

0o 00−1 × 10−5m/s

1−2 × 10−5m/s

2−3 × 10−5m/s

3−5 × 10−5m/s

>5 × 10−5m/s

M68/3 Jul 2006

26oW 24oW 22oW 20oW 18oW 16oW

16oN

18oN

20oN

22oN

(b)

(c)

(a)

(d)

Figure 3. Helium-derived upwelling velocities for cruise M91 off Peru (a) and for cruises M68/3, P347 and ATA3 off Mauritania (b–d).

Isobaths are drawn every 1000 m, and the area of the “coastal region” is shaded grey. Note that the uncertainty of the pointwise helium-derived

upwelling velocities is 68 % for the Peruvian and 100 % for the Mauritanian data (for details, see text).

10−5 ms−1 both for the Peruvian and for the Mauritanian re-

gions.

Off Peru, regions of strong coastal upwelling are found in

the northern area between 5 and 8◦ S. Another region with

high vertical velocities further south at 12–14◦ S is restricted

to the coast. Around 10◦ S and south of 15◦ S the upwelling

is weak or even vanishing.

The three cruises off Mauritania have a different regional

extension. Cruise P347 is restricted to the coast, but with

a very dense station spacing. The maximum of the upwelling

is located south of 18◦ N near the coast onshore of the 500 m

isobath. The other two cruises which cover a larger area

(M68/3 and ATA3) show that the upwelling near the coasts

persists up to 20◦ N near Cape Blanc. For these cruises, also

at some locations west of 18◦W enhanced upwelling is ob-

served. Combining the results from the three cruises from the

Mauritanian region, the whole picture is that in offshore di-

rection the vertical velocity first decreases and then increases

again at some locations.

Error estimation

The total error of the upwelling velocities is computed from

the error of the single terms in Eq. (1). This comprises the he-

lium values in boxes 1 and 2, C1, C2, the gas exchange and

the vertical mixing. The 3He / 4He ratio measured at the Bre-

men high-resolution mass spectrometer has a precision bet-

ter than 0.4 %. The standard deviation of duplicate samples

of the cruises presented here is even smaller, about 0.3 %.

This error of the 3He / 4He ratio is directly linked to the un-

certainty of C1, the 3He / 4He ratio in the mixed layer. The

number n1 of samples from the mixed layer for a single pro-

file, where C1 is derived from, varies between 1 and 3. So the

uncertainty of C1 is 0.3 %/√n1 ≈ 0.2 %.

Biogeosciences, 12, 7519–7533, 2015 www.biogeosciences.net/12/7519/2015/

R. Steinfeldt et al.: Coastal upwelling 7525

Below the mixed layer, again 1–3 helium measurements

leave its mark for the calculation of C2. Thus, the error of C2

is also about 0.2 %.

Another topic is the depth range that is chosen for the cal-

culation of C2. The concentration of 3He is typically increas-

ing with depth, so a deeper depth range for calculating C2

would lead to smaller upwelling velocities. This effect is par-

tially compensated by the diffusive term, as vertical mixing

below the mixed layer often decreases with depth, but in most

cases the effect of a depth increase of C2 will dominate. This

is in agreement with Kadko and Johns (2011), who postulate

a decrease of the upwelling velocity with depth. As noted

in Tanhua and Liu (2015), the upwelling velocities are valid

only for the depth where C2 is calculated, which is 5–25 m

below the mixed layer in our case.

The relative standard deviations of Kv for deep and shal-

low regions is about 100 %. Compared to this large value,

the error of the vertical helium-3 gradient can be neglected,

so the vertical mixing is estimated to vary by a factor of 2

from the calculated value.

For the gas exchange, an error of 30 % for the piston veloc-

ity is assumed. This error comprises both variations in wind

speed over the equilibrium timescale for the gas exchange

(see above) and errors in the parameterization of the piston

velocity. All these errors are listed in Table 1.

The influence of the errors of the input values on the up-

welling velocity w is non-linear (see Eq. 1). Thus the up-

welling velocities are calculated for adding and subtracting

the errors from the input values. In this way, a minimum and

a maximum upwelling velocity is computed, and the error is

assumed to be half of the difference. This is done for each

source of error separately. In the Peruvian upwelling, both

surface as well as subsurface helium-3 concentrations are

much higher than off Mauritania (Figs. 1 and 2). Thus, the

“signal to noise ratio” of C1 and C2 is larger for the region

off Peru, and the error estimation is done separately for both

upwelling regions.

The relative errors of w for each source of error are shown

in Table 1. Although the uncertainty of C1 and C2 is the

same, the resulting error in the upwelling velocity w is much

larger for C1 than for C2. This is because C1 appears in

Eq. (1) not only in the nominator (C1−C2), but also influ-

ences the magnitude of the gas exchange in the nominator of

Eq. (1). A change of C1 in one direction causes a change of

w in the same direction for both terms.

As was expected, the better signal to noise ratio of C1 and

C2 for the Peruvian area results in a smaller error for w com-

pared to the Mauritanian region. Also the uncertainty of w

resulting from the diapycnal mixing and the piston velocity

is smaller for the Peruvian region. Note that the uncertainty

of w due to vertical mixing is much smaller than the error of

the mixing itself, as in many cases the term for the vertical

diffusion in Eq. (1) is small compared to the gas exchange

term. The total error calculated as the quadratic sum of the

four error terms of the inferred upwelling velocity is 68 %

off Peru and 100 % off Mauritania.

5.2 Comparison between helium- and wind-derived

upwelling

We will now compare the upwelling velocities derived from

the helium method (wHelium) with those calculated directly

from the wind field (wWind). In the open ocean away from

coastal boundaries, the upwelling velocity at the base of the

Ekman layer can be computed directly from the wind stress

curl (see e.g. Gill, 1982)

w =1

ρ

(∂

∂x

(τy

f

)−∂

∂y

(τx

f

)), (4)

with water density ρ, Coriolis parameter f and the zonal and

meridional components of the wind stress τx and τy . Near the

coast, the lateral boundary is taken into account using a two-

layer model (Yoshida, 1955). The solutions for the velocity u

in the upper layer directed offshore and the vertical velocity

w have the form (Gill, 1982):

u=−τy

ρfH1

(1− e−x/a) (5)

w =τy

ρf ae−x/a . (6)

H1 is the depth of the upper layer (Ekman layer), and τydenotes the wind stress component parallel to the coast. The

spatial scale of the upwelling area is given by the first inter-

nal Rossby radius a. a is calculated for a two layer ocean

with densities ρ1 and ρ2. ρ1 is the density of the mixed layer.

For ρ2 we have chosen the mean density between the lower

boundary of the mixed layer and 500 m depth, which is ap-

proximately the lower boundary of the central water from

which the upwelled water originates. For each of the three

cruises, a mean value of a is calculated, as the stratification

and thus the Rossby radius between the cruises might differ.

The resulting values are a= 15 km for M91, a= 16 km for

M68/3, a= 10 km for P347 and a= 12 km for ATA3. The

magnitude of the wind-driven coastal upwelling velocities

at each station depends on the choice of the Rossby radius

a in Eq. (6). The total vertical transport integrated over the

coastal area, i.e. a distance of several Rossby radii, is al-

most independent from a. The alongshore velocity (coastal

jet) in Eq. (5) increases linearly with time, so this solution

does not represent a steady state. More complex solutions

can be found (Fennel, 1999) where the increase of the coastal

jet is limited due to the generation of coastal trapped Kelvin

waves.

x in Eqs. (5) and (6) denotes the distance from the coast.

The continental shelf in the study area is relatively broad, and

the mixed layer in this region has a mean depth of 10–25 m.

The upwelled water has to be supplied from below the mixed

layer, but this is only possible if the water depth is consider-

ably larger than the mixed layer (more correct: Ekman layer)

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7526 R. Steinfeldt et al.: Coastal upwelling

depth itself. We thus assume a minimum water depth of 50 m

(two times the mixed layer depth) for wind-driven upwelling

to occur and set x as the distance from the 50 m isobath. In

this case, wWind calculated according to Eq. (6) is compa-

rable in magnitude with wHelium (Fig. 4a). Setting x = 0 di-

rectly at the coast leads to wWind being 1 order of magnitude

smaller. As wind-driven vertical velocity wwind which has to

be compared with wHelium we choose the maximum value of

both calculations according to Eqs. (4) and (6), which means

that either the influence of the coastal boundary is dominat-

ing the vertical Ekman velocity or the wind stress curl over

the open ocean.

wWind is calculated from the gridded daily wind data

which are interpolated onto the station locations in the same

way as for calculating the gas exchange velocity. Also for

wWind the temporal mean over the period that is given by

the gas exchange timescale (n days, see Sect. 3) is taken.

So wWind covers the same timescale of n days as wHelium.

Nevertheless, there is a difference in the interpretation of

the temporal mean of wWind and wHelium. Whereas wHelium

has to be interpreted as the Lagrangian mean following the

patch of surface water, wWind is the Eulerian mean. An-

other difference between wWind and wHelium is the fact that

wHelium strictly speaking describes the entrainment velocity:

wHelium = wWind− ∂zmld/∂t . Here, zmld denotes the lower

boundary of the mixed layer.

To overcome the difference between these methodological

differences and also to reduce the large error of the pointwise

wHelium data, regional mean values of wWind and wHelium are

calculated. For these mean values, the differences between

Eulerian and Lagrangian mean as well as between vertical

and entrainment velocity should be reduced. The regions are

the area within the 50 km distance to the 50 m isobath, where

the boundary solution from Eq. (6) is typically larger than

the open ocean result from Eq. (4), all stations with larger

distance from the coast make up the other region. These areas

will be referred to as “coastal”, and “offshore” respectively.

The error of the regional mean values of wWind and wHelium

is the standard deviation of the pointwise velocity data within

the region with helium data divided by the square-root of the

number of those stations. For the helium-derived upwelling,

this error is typically larger than dividing the 100 % error of

the single measurements by the square root of the number

of data points. The reason is that vertical velocities in one

region are varying by up to an order of magnitude.

All mean values for the coastal and offshore regions for

each cruise are given in Table 2 and represented graphi-

cally in Fig. 4. For the coastal regions, mean upwelling

velocities are of the order of 10−5 ms−1, whereas for the

offshore regions they vary between 10−5 ms−1 and only

10−6 ms−1. The coastal values for wHelium and wWind off

Mauritania agree for all three cruises within their error bars.

The winter cruise P347 shows the smallest coastal upwelling

1.4± 0.4× 10−5 m s−1. The other winter cruise ATA3 and

also the summer cruise M68/3 have larger values 2.1±0.8×

Upw

ellin

g V

eloc

ity [1

0−5 m

/s]

Hel

ium

Win

d

Hel

ium

red

. Hel

ium

Win

d

Hel

ium

Win

d

Hel

ium

Win

d

0

1

2

3

4

5

6 M91 Dec 2012M68/3 Jul 2006P347 Jan 2007ATA3 Feb 2008

Upw

ellin

g V

eloc

ity [1

0−5 m

/s]

Hel

ium

Win

d

Hel

ium

Win

d

Hel

ium

Win

d

Hel

ium

Win

d

0

0.5

1

1.5

2

2.5M91 Dec 2012M68/3 Jul 2006P347 Jan 2007ATA3 Feb 2008

(a)

(b)

Figure 4. Mean values of helium- and wind-derived upwelling ve-

locities for the coastal (a) and offshore (b) areas of cruises M91,

M68/3, P347 and ATA3. The vertical black line indicates the stan-

dard deviation of the mean. “Helium red” in (a) means that the

helium-derived upwelling is calculated with reduced gas exchange

velocity due to the presence of surface organic films, for details see

text.

10−5–2.4±1.5×10−5 ms−1. These numbers, however, agree

within their error bars. This implies that no seasonal variation

of the upwelling has been observed, but the limited number

of three cruises might not be representative for the respec-

tive season. Offshore, only for cruise P347 wHelium is small

(0.2× 10−5 ms−1). For all Mauritanian cruises wHelium sur-

passes wWind in the offshore region.

For the Peruvian area, the differences between wHelium

and wWind show an opposite behaviour: they are relatively

small for the offshore region, but large for the coastal area (≈

1.1×10−5 ms−1 for wWind in contrast to≈ 2.7×10−5 ms−1

for wHelium). One reason for the high value of wHelium might

be an overestimation of the gas exchange velocity in Eq. (3)

due to the presence of organic surface films (surfactants).

These films have been observed on cruise M91, and their

damping behaviour on surface waves have been shown in

Kiefhaber et al. (2015). There, the observed mean square

slope of the waves was found to be overestimated by the

parameterization for clean water. Most measurements are

in the range spanned by clean water and slick parameteri-

zations. Surfactants also drastically reduce the transport of

gases across the water surface. A parameterization of the gas

transfer velocity in the presence of surfactant is given in Tsai

and Liu (2003). However, there the gas transfer for the case

with and without surface films are based on the parameteri-

Biogeosciences, 12, 7519–7533, 2015 www.biogeosciences.net/12/7519/2015/

R. Steinfeldt et al.: Coastal upwelling 7527

Table 2. Mean upwelling velocities w.

M91 M68/3 P347 ATA3

coastal coastal surf. offshore coastal offshore coastal offshore coastal offshore

] prof. w > 0 14 12 12 5 8 17 4 2 7

] prof. w ≤ 0 5 7 18 3 9 11 6 1 1

wHelium [10−5 ms−1] 2.7± 0.6 1.0± 0.3 0.6± 0.2 2.1± 0.7 0.7± 0.3 1.4± 0.4 0.2± 0.1 2.4± 1.5 1.4± 0.9

wWind [10−5 ms−1] 1.3± 0.2 – 0.4± 0.1 3.7± 1.7 0.3± 0.1 1.2± 0.3 0.02± 0.01 1.6± 0.9 0.07± 0.04

zation form Liss and Merlivat (1986). For low wind speeds

as have been prevailing during cruise M91, this formula re-

sults in smaller gas transfer velocities than Eq. (3) even for

the normal case without surfactants. Including the reduction

factor r from Tsai and Liu (2003) (r = 0.56 U−0.1310 ) for the

case with surfactant, the resulting gas exchange and thus the

helium-derived vertical velocities would almost vanish. We

thus adopt the reduction factor from Tsai and Liu (2003), but

apply it to the gas transfer velocity from Nightingale et al.

(2000), Eq. (3). The resulting mean upwelling velocity for

the coastal area is also given in Table 2 and Fig. 4a and is in

good agreement with the wind-derived value wWind. Figure 5

shows the distribution wHelium for the case of the reduced

gas transfer velocity. Comparison with Fig. 3a for the stan-

dard gas exchange shows that the overall pattern of the dis-

tribution of the vertical velocities remains unchanged, only

the coastal values are smaller. In Kiefhaber et al. (2015) not

all data points are influenced by surfactants. As illustrated in

Fig. 4a, the helium-derived upwelling velocity based on re-

duced gas exchange at all coastal stations is slightly smaller

than the wind-derived one, indicating that the effect of sur-

factants is overestimated when being applied to all stations.

One could argue that the enhanced δ3He values in the off-

shore region are the remnants from helium-3 rich water origi-

nating in the coastal upwelling and then have being advected

offshore. Taking into account the equilibrium timescale for

the helium gas exchange, after about 20 days the mixed-layer

disequilibrium of helium-3 should have decreased to about

10 % of the value from the upwelled water (< 0.1 % for the

Mauritanian and < 0.2 % for the Peruvian upwelling). As-

suming an advection velocity of ≈ 10 cms−1 in offshore di-

rection, the water could move about 200 km away from the

coast over the 20-day time period. For the Peruvian area,

most stations are within this distance, so an influence from

the coastal upwelling on the offshore helium-3 values can-

not be excluded. For the Mauritanian region, enhanced δ3He

values can be found even west of 18◦W, too far west to be

remnants from the coastal upwelling. Possible explanations

for the large offshore vertical velocities will be given below.

5.3 Other upwelling mechanisms

The discrepancies between the wind- and helium-derived up-

welling velocities especially in the offshore region suggest

the existence of additional upwelling mechanisms. These

M91 Dec 2012

85oW 80oW 75oW 70oW 20oS

15oS

10oS

5oS

0o 00−1 × 10−5m/s

1−2 × 10−5m/s

2−3 × 10−5m/s

3−5 × 10−5m/s

Figure 5. Helium-derived vertical velocity for cruise M91. As in

Fig. 3a, but applying a reduced gas transfer velocity in the coastal

area due to the presence of surface films.

would be included in the helium-derived vertical velocities,

but not in the purely wind-driven ones. For the region at

the Mauritanian coast, large helium-derived velocities for

cruises M68/3 and ATA3 are located south of Cape Blanc

near 20◦ N (Fig. 3b and d). These maxima do not show corre-

sponding high values in the wind-driven upwelling. Mazzini

and Barth (2013) concluded from a model study that flow–

topography interaction is upwelling favourable downstream

of capes, a situation which is given south of Cape Blanc fol-

lowing the south westward direction of the Canary Current.

Thus in this case the wind-derived vertical velocity might be

an underestimation.

Another explanation for the high vertical velocities from

the helium method which surpass the Ekman derived values

by up to 1 order of magnitude is eddy-induced upwelling.

Several mechanisms for this phenomenon are possible: up-

lift of isopycnals close to the mixed layer, as they occur in

cyclones and mode water eddies. Stramma et al. (2013) ob-

served subsurface chlorophyll maxima in such types of ed-

dies in the Peruvian upwelling on cruise M90, just 1 month

prior to the cruise M91 discussed here, and ascribed them to

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7528 R. Steinfeldt et al.: Coastal upwelling

Sea level anomaly [cm]

wH

e − w

Win

d [10−

5 m/s

]

−5 0 5 10 15−4

−2

0

2

4

6

8M91 Dec 2012M68/3 Jul 2006P347 Jan 2007ATA3 Feb 2008

Figure 6. Difference between helium- and wind-derived vertical ve-

locity against sea level anomaly at each station for the coastal (filled

circles) and offshore (open circles) region of cruises M91, M68/3,

P347 and ATA3. Note that the uncertainty of the helium-derived up-

welling velocities is 68 % for the Peruvian and 100 % for the Mau-

ritanian data.

this mechanism. The helium method is however not able to

catch this process as long as the uplifting of isopycnals does

not lead to an intrusion of subsurface water into the mixed

layer. In the opposite direction, when the upwelled water

leaves the coastal area, the mixed layer might deepen, leading

to entrainment of water from below. This process would be

classified as upwelling by the helium method. However, no

correlation between enhanced offshore upwelling and deep

mixed layers can be found from the data off Mauritania.

Ekman suction due to wind stress on the eddy is also

suggested to foster upwelling (Martin and Richards, 2001).

Here, at the side of the eddy where the eddy flow is in

the same direction as the wind, the wind stress is reduced,

whereas it is enhanced on the opposite side of the eddy,

where wind and eddy flow are in the opposite direction. This

difference in wind stress between both sides of the eddy in-

duce an upwelling for anticyclones and mode water eddies

and a downwelling for cyclones. Another mechanisms also

described in Martin and Richards (2001) is upwelling by

ageostrophic circulation resulting from a perturbation of the

eddy flow field.

In order to investigate a possible influence of eddies on

the helium-derived upwelling velocity, the gridded AVISO

sea level anomaly (SLA) has been interpolated on the sta-

tion location. Figure 6 shows the interpolated SLA against

1w = wHelium−wWind, i.e. the portion of the vertical ve-

locity, which is not explained by wind-driven upwelling. For

the coastal data, no relation between SLA and 1w can be

found. For the offshore data, towards the centre of anticy-

clones (SLA> 5 cm), the upwelling is slightly reduced. At

offshore points with smaller SLA (5cm> SLA>−5 cm),

1w is either positive or close to zero. If the two data points

with wHelium > 4× 10−5 ms−1 are regarded as outliers, the

offshore data form a triangular pattern with the edges at

[−4cm/0 ms−1], [1 cm/3× 10−5 ms−1] and [5 cm/0 ms−1].

The eddy–wind interaction can thus be ruled out as mech-

anisms responsible for the upwelling, as this only works for

positive SLA. Another reason is that even in that case, for

the moderate wind speeds observed during all cruises (u10 <

10 ms−1), the resulting upwelling would only be of the or-

der of 10−6 ms−1. A strong upwelling of order 10−5 ms−1

as observed by McGillicuddy Jr. et al. (2007) only occurs for

high windspeeds (u10 > 10 ms−1).

The spatial distribution of SLA and 1w at the offshore

data points is shown in Fig. 7 for each cruise. Whereas in

Fig. 6 weekly means are used for SLA, in Fig. 7 the mean

from the weekly SLA over the duration of the cruises is

shown in order to get only one map per cruise and thus

a comprehensive picture. Figure 7 illustrates that enhanced

upwelling occurs at the edge of eddies at locations of small

SLA. These enhanced vertical velocities are, however, not

a general feature of eddy boundaries. We conclude that

ageostrophic instabilities that might occur at the edge of ed-

dies are the main mechanism for eddy-induced upwelling.

On the other hand, towards the centre of anticyclones up-

welling is weakened.

5.4 Nutrient fluxes

The coastal regions off Peru and Mauritania belong to

the most productive areas of the world ocean. We thus

consider the relation between nutrient supply into the

mixed layer from vertical transports (both advective and

diffusive) and net primary production (NPP) observed

from satellites (http://www.science.oregonstate.edu/ocean.

productivity/index.php). The advective and diffusive phos-

phate fluxes into the mixed layer are computed in the same

way as the helium-3 flux: the concentration of phosphate in

box 2 (C2), the phosphate gradient and the vertical diffusivity

are calculated as vertical mean over the depth range 5–25 m

below the mixed layer. If no microstructure data are avail-

able, the diffusion coefficient is set to the regional mean (see

Sect. 3). As advective velocities in the Peruvian upwelling

wHelium derived from the reduced gas exchange is used, these

values are in better agreement with wWind.

Similar to the SLA data, the 8 day mean values of satellite-

derived productivity are interpolated on the station locations.

In order to allow for a reaction of the productivity to changes

in nutrient supply, the productivity data are shifted in time

by half a week. For a quantitative comparison between NPP

and phosphate fluxes, the latter are converted to carbon units

by multiplying with the Redfield ratio of 117 from Anderson

and Sarmiento (1994).

The spatial distribution of NPP and vertical carbon trans-

port is shown in Fig. 8 for each cruise. As for SLA, the NPP

values over the time period of each cruise are averaged to get

one comprehensive map per cruise. Here, at least a qualitative

correlation between NPP and carbon flux can be observed. At

the northern end of the Peruvian area, e.g. enhanced carbon

fluxes reach from the onshore up to the offshore end of the

Biogeosciences, 12, 7519–7533, 2015 www.biogeosciences.net/12/7519/2015/

R. Steinfeldt et al.: Coastal upwelling 7529

P347 Jan 2007

26oW 24oW 22oW 20oW 18oW 16oW

16oN

18oN

20oN

22oN

M68/3 Jul 2006

26oW 24oW 22oW 20oW 18oW 16oW

16oN

18oN

20oN

22oN

M91 Dec 2012

85oW 80oW 75oW 70oW 20oS

15oS

10oS

5oS

0o <=00−2 × 10−5m/s

>4 × 10−5m/s

SLA[cm]

−201512.5−10−8−6−4−2024681012.51520

ATA3 Feb 2008

26oW 24oW 22oW 20oW 18oW 16oW

16oN

18oN

20oN

22oN

(a)

(c)

(b)

(d)

Figure 7. Mean sea level anomaly over the time period of the respective cruise for M91 (a), M68/3 (b), P347 (c), and ATA3 (d). The

difference between helium- and wind-derived vertical velocity is indicated by the grey dots. For details see text.

sections, and the area with enhanced NPP also stretches rela-

tively far offshore in this northern area. For cruise M68/3, the

offshore area with striking high NPP around 20◦W might be

fostered by the relatively large vertical nutrient transport ob-

served at the two stations near 19◦W. The reason for the high

nutrient fluxes at this location is the above-mentioned eddy-

induced upwelling. The stations near 18◦W, 16.5◦ N are at

least in close vicinity to the area with high NPP along the

coast. For cruise P347 the calculated nutrient–carbon fluxes

reflect the decrease of NPP in offshore direction. The limited

number of stations with nutrient fluxes for ATA3 hamper to

find a correlation with the underlying NPP field.

The spatial misfit between NPP and vertical nutrient sup-

ply might also be due to the temporal delay between them.

Over this delay time, the upwelled water is advected hori-

zontally, so NPP and nutrient flux are not expected to ap-

pear exactly at the same location. The lack of correlation be-

tween upwelling (local forcing) and primary production has

also been found in a study by Carr and Kearns (2003). They

analysed the governing factors for the biological production

in eastern boundary current systems and found even negative

correlations between local forcing and primary production

both for the northern part of the Humboldt Current off Peru

(5–15◦ S) and the southern part of the Canary Current off

Mauritania (11◦ S–20◦ N) (Carr and Kearns, 2003, Table 3).

The important parameter for primary production is the nu-

trient flux at the base of the euphotic zone. However, neither

the depth of the euphotic zone nor the vertical velocity w at

this location is precisely known. Typically, the euphotic zone

is deeper than the relatively shallow mixed layers observed in

the upwelling regions. Assuming a decrease of w with depth

and an increase of nutrient concentrations, these two gradi-

ents are counteracting in their effect on the nutrient flux, so

the numbers presented here are considered as approximations

of the nutrient flux in the euphotic zone.

Values for NCP in the Peruvian and Mauritanian up-

welling have been calculated by Minas et al. (1986),

0.59 gCm−2 d−1 off Peru and 2.51 gCm−2 d−1 off Mauri-

tania. For comparison, the regional mean values of the ver-

tical carbon fluxes from this study are given in Table 3, di-

vided into a diffusive and an advective part. As in Minas et al.

(1986), the value for the Peruvian upwelling is smaller than

for Mauritania (1.3 vs. 1.6–2.1 gCm−2 d−1), but the differ-

ence is much less pronounced.

The nutrient fluxes into the mixed layer by vertical ad-

vection and vertical mixing are of similar magnitude for

the Peruvian offshore area and in most cases both for the

coastal and offshore region off Mauritania. Over the Mauri-

tanian shelf break, the vertical diffusivity is largely enhanced

due to tide–topography interactions (Schafstall et al., 2010),

which explains the high diffusive fluxes. Our result of 0.8–

www.biogeosciences.net/12/7519/2015/ Biogeosciences, 12, 7519–7533, 2015

7530 R. Steinfeldt et al.: Coastal upwelling

M91 Dec 2012

85oW 80oW 75oW 70oW 20oS

15oS

10oS

5oS

0o <0.50.5−11−2

NPP[g C−1m−2d−1]

00.20.40.60.811.21.41.61.82345715 M68/3 Jul 2006

26oW 24oW 22oW 20oW 18oW 16oW

16oN

18oN

20oN

22oN

P347 Jan 2007

26oW 24oW 22oW 20oW 18oW 16oW

16oN

18oN

20oN

22oN

ATA3 Feb 2008

26oW 24oW 22oW 20oW 18oW 16oW

16oN

18oN

20oN

22oN

(a)

(b)

(d)(c)

Figure 8. Mean net primary production over the time period of the respective cruise for M91 (a), M68/3 (b), P347 (c), and ATA3 (d). The

vertical phosphate transport (advective plus diffusive part) into the mixed layer is converted to carbon units by the Redfield ratio and indicated

by the grey dots. Note that the error of these fluxes is 100 %. For details see text.

Table 3. Mean advective, diffusive, and total PO4 fluxes into the mixed layer converted to carbon units [g C m−2 d−1] via the Redfield ratio.

M91 M68/3 P347 ATA3

coastal offshore coastal offshore coastal offshore coastal offshore

adv. 1.0± 0.3 0.4± 0.1 1.1± 0.4 0.7± 0.3 0.8± 0.3 0.2± 0.1 – 0.4± 0.3

diff. 0.3± 0.1 0.5± 0.1 0.9± 0.2 0.9± 0.2 0.8± 0.4 0.2± 0.1 0.9± 0.3 0.08± 0.03

total 1.3± 0.3 0.9± 0.1 2.1± 0.5 1.6± 0.4 1.6± 0.5 0.4± 0.1 – 0.5± 0.3

For cruise ATA3, no advective nutrient flux for the coastal area has been computed. Only one station with helium and phosphate data for that regions is

available, which is not sufficient to calculate a mean value.

0.9 gCm−2 d−1 is almost identical with the study in Schaf-

stall et al. (2010) for the same region, but not exactly the

same subset of cruises. Further offshore, the vertical diffu-

sivity drops by more than 1 order of magnitude, which is

not compensated by the slightly larger subsurface increase

of nutrients in the offshore areas (Fig. 2d–f). This is differ-

ent for the Peruvian region, where coastal and offshore dif-

fusivities reach the same magnitude at some stations. Here,

the diffusive flux offshore is even higher than near the coast,

and the coastal flux is dominated by the advective contribu-

tion. According to Carr and Kearns (2003), the open ocean

productivity is reduced by a factor of 10 compared to the

coastal upwelling areas. Such small values of vertical nu-

trient transport have been found in Haskell II et al. (2015)

further offshore off Peru (0.01–0.1 gCm−2 d−1). This indi-

cates the transition from the upwelling regime towards the

oligotrophic subtropical gyre. In the offshore regions adja-

cent to the coastal boundary analysed here, the vertical nu-

trient transport is also smaller than at the coast, but only by

a factor of 1.3 (Peru, cruise M91) to 4 (Mauritania, cruise

P347). The offshore filaments of enhanced productivity ob-

served from satellite (Fig. 8) might thus not only be fed by

horizontal advection of nutrients out of the coastal zone, but

also by nutrient input from below the mixed layer.

The general importance of eddies for the nutrient supply

has been shown by Oschlies and Garçon (1998) in a model

Biogeosciences, 12, 7519–7533, 2015 www.biogeosciences.net/12/7519/2015/

R. Steinfeldt et al.: Coastal upwelling 7531

for the Atlantic. In another model study Gruber et al. (2011)

show that the presence of eddies reduces the nutrient trans-

port into the euphotic zone and thus the productivity in

eastern boundary upwelling systems compared to the non-

eddying case in the region adjacent to the coast (0–500 km).

Further offshore (500–1000 km), however, the eddy field in-

duces an increase in primary production and carbon export.

6 Conclusions

Vertical velocities w for the eastern boundary current up-

welling systems off Peru and Mauritania regions have been

determined by using the 3He / 4He disequilibrium in surface

waters. The mean upwelling velocity over the coastal regions

varies between 1.1± 0.3× 10−5 and 2.4± 1.5× 10−5 ms−1

and is similar for both regions. In the equatorial Atlantic

Rhein et al. (2010) and Kadko and Johns (2011) found verti-

cal velocities reaching from 0.6× 10−5 to 2.6× 10−5 ms−1.

Thus both the equatorial and the coastal upwelling are of sim-

ilar strength.

For the Mauritanian area, Tanhua and Liu (2015) have

calculated upwelling velocities using the same box model

as here, but CFC-12 and SF6 as tracers. For a winter and

spring cruise, the vertical velocities are between 0 and 11×

10−5 ms−1, comparable with our results. For the summer

cruise M68/3, Tanhua and Liu (2015) find no indications for

upwelling. Their data set, however, only contains a few sta-

tions along 18◦ N. Our coastal stations along that line also

have low vertical velocities (between 0 and 2× 10−5 ms−1),

whereas directly north and around 20◦ N the helium-derived

upwelling exceeds 3× 10−5 ms−1.

Haskell II et al. (2015) investigate upwelling in the East-

ern Tropical South Pacific at a few stations based on the7Be method. Their vertical velocities are also in the order of

10−5 ms−1 (0–3×10−5 ms−1), whereby the maximum value

of 3×10−5 ms−1 might be due to “anomalous 7Be measure-

ments”.

An independent estimate of the upwelling velocity can be

inferred by Ekman theory (Gill, 1982). Near the coast, the

agreement between wind- and helium-derived mean vertical

velocities is fair, if a “minimum” water depth of 50 m for up-

welling to occur is chosen. For the Peruvian upwelling, the

vertical velocity-derived from the helium method in conjunc-

tion with the gas exchange parameterization of Nightingale

et al. (2000) leads to an overestimation of the upwelling, as

the gas exchange is reduced due to the presence of surface

surfactants (Kiefhaber et al., 2015). Assuming a reduction of

the piston velocity in the whole coastal area off Peru leads

to a small underestimation of the upwelling, but helium- and

wind-derived values still agree within their errors.

In contrast to the broad agreement between the mean

upwelling derived from the helium and the wind method,

at greater distances from the Mauritanian coast large dis-

crepancies occur. Here, the helium-derived upwelling still

reaches 1× 10−5 ms−1, whereas the wind-driven upwelling

from Ekman suction is smaller by up to one order of mag-

nitude. Haskell II et al. (2015) find an upward velocity of

1× 10−5 ms−1 at 10◦ S, 100◦W, which is another indica-

tion for enhanced upwelling in the open ocean. One possible

mechanism is eddy-induced upwelling by perturbations of

the eddy flow field (Martin and Richards, 2001). This view is

supported by sea level anomalies, which are moderate at the

locations of maximum upwelling, i.e. the largest upwelling

is found at the edge of eddies and not in their centre.

Vertical advection and diapycnal mixing are the most im-

portant mechanisms for the transport of substances from

the interior ocean into the surface mixed layer. As a con-

sequence, the upwelling regimes belong to the most pro-

ductive ocean regions. Due to the high vertical diffusivities

at the shelf break off Mauritania (Schafstall et al., 2010),

in that area the diffusive nutrient flux is of the same mag-

nitude as the advective one. Both types of fluxes together

are equivalent to a carbon flux of 1.3± 0.3 gCm−2 d−1 for

the Peruvian and 1.6–2.1± 0.5 gCm−2 d−1 for the Mauri-

tanian region. The upwelling also leads to enhanced nutri-

ent fluxes in the offshore region, which can reach up to

1.5±0.4 gCm−2 d−1 off Mauritania during the cruise M68/3

and 1.0± 0.1 gCm−2 d−1 off Peru for cruise M91.

Acknowledgements. This work is part of the German research

project Surface Ocean Processes in the Anthropocene (SO-

PRAN), funded by the German Federal Ministry of Education

and Research, BMBF, grants 03F0462D, 03F0611D, 03F0662D

(M. Rhein, R. Steinfeldt, J. Sültenfuß) and grants 03F0462A,

03F0611A, 03F0662A (M. Dengler, T. Fischer). The authors

thank captain and crews of the cruises M91, M68/3, P347,

and ATA3 for their professional field support. The wind speed

data are derived from ftp://ftp.ifremer.fr/ifremer/cersat/products/

gridded/mwf-ascat/data/daily and ftp://ftp.ifremer.fr/ifremer/

cersat/products/gridded/mwf-quikscat/data/daily for the time

period of the Peruvian and Mauritanian cruises respectively.

The satellite data of primary production can be found at

http://www.science.oregonstate.edu/ocean.productivity/index.php.

The altimeter products were produced by Ssalto/Duacs

and distributed by Aviso with support from CLS-Cnes

(http://www.aviso.oceanobs.com/duacs/). The contribution

from the group of Arne Körtzinger (GEOMAR), making the

nutrient data available, is gratefully acknowledged. We also thank

two anonymous reviewers for their valuable comments.

The article processing charges for this open-access

publication were covered by the University of Bremen.

Edited by: H. Bange

www.biogeosciences.net/12/7519/2015/ Biogeosciences, 12, 7519–7533, 2015

7532 R. Steinfeldt et al.: Coastal upwelling

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